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dynamics during the ESCOMPTE campaign
F. Saïd, Aurore Brut, Bernard Campistron, F. Cousin
To cite this version:
F. Saïd, Aurore Brut, Bernard Campistron, F. Cousin. Investigation of the atmospheric boundary layer
dynamics during the ESCOMPTE campaign. Annales Geophysicae, European Geosciences Union,
2007, 25 (3), pp.597-622. �hal-00330120�
www.ann-geophys.net/25/597/2007/
© European Geosciences Union 2007
Annales Geophysicae
Investigation of the atmospheric boundary layer dynamics during the ESCOMPTE campaign
F. Sa¨ıd1, A. Brut2, B. Campistron1, and F. Cousin3
1
Centre de Recherches Atmosph´eriques, UMR CNRS 5560, 8 route de Lannemezan, 65300 Campistrous, France
2
Centre d’Etudes Spatiales de la BIOsph`ere UMR 5126, 18 avenue Edouard Belin, 31401 Toulouse Cedex 9, France
3
Laboratoire d’A´erologie, UMR CNRS 5560, 14 avenue Edouard Belin, 31400 Toulouse, France
Received: 25 April 2005 – Revised: 7 February 2007 – Accepted: 27 February 2007 – Published: 29 March 2007
Abstract. This paper presents some results about the behav-
ior of the atmospheric boundary layer observed during the ESCOMPTE experiment. This campaign, which took place in south-eastern France during summer 2001, was aimed at improving our understanding of pollution episodes in rela- tion to the dynamics of the lower troposphere. Using a large data set, as well as a simulation from the mesoscale non- hydrostatic model Meso-NH, we describe and analyze the atmospheric boundary layer (ABL) development during two specific meteorological conditions of the second Intensive Observation Period (IOP). The first situation (IOP2a, from 22 June to 23 June) corresponds to moderate, dry and cold northerly winds (end of Mistral event), coupled with a sea- breeze in the lower layer, whereas sea-breeze events with weak southerly winds occurred during the second part of the period (IOP2b, from 24 June to 26 June).
In this study, we first focus on the validation of the model outputs with a thorough comparison of the Meso-NH simula- tions with fields measurements on three days of the IOP: 22 June, 23 June and 25 June. We also investigate the structure of the boundary layer on IOP2a when the Mistral is super- imposed on a sea breeze. Then, we describe the spatial and diurnal variability of the ABL depths over the ESCOMPTE domain during the whole IOP. This step is essential if one wants to know the depth of the layer where the pollutants can be diluted or accumulated. Eventually, this study intends to describe the ABL variability in relation to local or mesoscale dynamics and/or induced topographic effects, in order to ex- plain pollution transport processes in the low troposphere.
Keywords. Atmospheric composition and structure (Pol-
lution – urban arid regional; Troposphere – composition and chemistry) – meteorology and atmospheric dynamics (Mesoscale meteorology)
Correspondence to: A. Brut
(aurore.brut@cesbio.cnes.fr)
1 Introduction
Prediction of pollution events and implementation of adapted air quality policies require a considerable improvement in the understanding of chemical and dynamical processes which lead to high pollutant concentrations in the atmosphere. Pe- culiar attention must be paid to the processes which take place in coastal zones, where cities and industries often con- centrate. In addition, coastal regions are also subject to wind circulations that are both synoptically induced and local- ized, or diurnal and terrain-induced (Rosenthal et al., 2003;
Angevine et al., 2004).
The Mediterranean basin in Europe is surrounded by high coastal mountains while it is heavily urbanized and industri- alized. The meteorological processes in summer are dom- inated by the Azores anticyclone, situated in south-western Europe and characterized by warm temperatures and moder- ate winds favorable to the onset of sea-breeze circulations.
Several European projects have been conducted around the Mediterranean basin to document the circulation of air pol- lutants: among them are the MEso-meteorological Cycles of Air Pollution in the Iberian Peninsula, MECAPIP, 1988–
1991 (Mill`an et al., 1996, 2000), the REgional Cycles of Air Pollution in the west-central Mediterranean Area, RECAPMA, 1990–1992 (Gangoiti et al., 2001) and the South European Cycles of Air Pollution, SECAP, 1992–1995 (Mill`an et al., 1997). These experimental projects showed that during daytime and summer conditions, the combina- tion of sea breezes with up-slope winds, due to the strong heating on the mountains, may inject pollutants up to the low (2–3 km) or even middle troposphere (3–5 km), form- ing an accumulation of ozone layers over the continent.
This aged pollution travels at upper levels and may come
back down to the sea or to the coastal areas, through sub-
sidence processes, such as the compensatory return flow of
the up-slope and sea-breeze circulations. These scenarios are
likely to occur in the area selected for the European cam-
paign ESCOMPTE “Exp´erience sur Site pour CONtraindre
les Mod`eles de Pollution atmosph´erique et de Transport d’Emissions” (Cros et al., 2004), which took place between the beginning of June 2001 and mid-July 2001. The speci- ficity of this project, in comparison with the previous ones, was to reinforce the experimental as well as the numerical frame. The main objective of the ESCOMPTE experiment was indeed to establish a detailed three-dimensional database of primary pollutant emissions, together with the dynamics and chemical composition of the atmosphere, in order to validate and improve numerical modeling of chemistry and transport. A 120
×120 km area centered on the coastal city of Marseille and the Berre pond, one of the most polluted region in France, was selected to host the experiment. This region presents high occurrences of photochemical pollution events due to its dense industrialization and urbanization, combined with the hot and sunny weather conditions of summertime.
The transport of the plumes is also highly influenced by the characteristics of topography and subject to secondary cir- culation systems, such as land/sea-breeze, slope winds and valley-mountain winds. Thus, the second Intensive Observa- tion Period (IOP2) of the ESCOMPTE campaign that lasted from 22 June to 26 June (hereafter identified as J22 to J26) was selected for the present study since the level of pollution was the highest of the experiment. During this period, a large amount of aircraft, radars and radio-soundings measurements was available. Using this large data set, in addition to numer- ical simulations, we intended to investigate the meteorolog- ical processes which affect the atmospheric concentration of pollutants (accumulation, dispersion and transport). These features include the development and thickness of the Atmo- spheric Boundary Layer (hereafter, ABL) and the effect of mesoscale circulation. Our analysis of atmospheric dynam- ics is mainly focused on two kinds of typical meteorological conditions that occurred in south-eastern France: just before J22, a strong northerly regional wind called Mistral had been blowing in the Rhone valley. It weakened on J22 and J23 which enabled the sea breeze to impact the lower layers near the coast. Gu´enard et al. (2005) showed that during the day- time, the coastal Mistral was affected by the sea breeze and by thermally driven convection which could lift up the Mis- tral above 2 km. This is consistent with our results, although we observe that the boundary layer development is rather dif- ferent over the region for these two days.
After that, the meteorological situation turned to weak southerly winds with a diurnal sea-breeze circulation from J24 to J26. For these typical conditions, differences appeared on the spatial and temporal variability of the ABL, therefore influencing the possibilities of exchanges between the ABL and the free troposphere.
The present paper is organized as follows: the instrumental set-up and data processing are presented in Sect. 2. Section 3 proposes a brief description of the meteorological conditions observed during the two periods of IOP2. In Sect. 4, simula- tion and experimental results of IOP2 are compared. It is also an opportunity to point out different behaviors of the ABL
development between J22 and J23 (both days correspond- ing to a situation of Mistral wind combined with a weak sea breeze near the coast) and to detail a pure sea-breeze case (J25). Section 5 addresses a point of major interest which concerns the variability of the depth of the ABL, since it im- pacts the layer where pollutants are likely to dilute or accu- mulate. Finally, we describe in Sect. 6 the resulting fields of ozone concentration and comment on them in light of the previous results.
2 Data and tools
The wide instrumental set-up deployed over the region was designed to document dynamics as well as chemistry, in- cluding aerosols. For the present work, we used meteoro- logical and chemical measurements collected on board three airplanes. In addition, three-dimensional fields of the wind components were provided by a profiler network of UHF radars, and vertical profiles of meteorological parameters were obtained from radio-soundings launched from several places over the area. The location of the various instrumen- tal sites is shown on Fig. 1.
The complex topography of the region has to be noticed.
As the region is situated at the foothill of the Alps (a large and high ridge exceeding 3000 m), several mountainous mas- sifs whose altitudes can reach 1000 m are encountered in the eastern and northern parts of the domain. To the west, the Crau plain and the wide Rhˆone valley (north-south orienta- tion) correspond to flat and low terrain, except for the 450- m high Alpilles ridge which crosses the Rhˆone valley. To the north, the Apt and Durance valleys are narrow channels that sink into the mountains: the Durance valley is first ori- ented north-east/south-west (Vinon-Cadarache-Pertuis) and then bends towards a west-east direction, parallel to the Apt valley, but on the other side of the Lub´eron massif. In ad- dition to this complex topography, the presence of the sea to the south is another factor which will play a role in local circulations.
2.1 UHF wind-profilers
The wind profilers network is made up of four Ultra High Frequency (UHF) radars (DEGREWIND PCL1300) located in Saint Chamas, Marignane, Aix-les-Milles and Marseille.
The profilers have the same technical features: a 4-kW peak-
power, a 1238-MHz transmitted frequency and a 150-m pulse
width. Except for the profiler installed at Aix-les-Milles,
which only used three beams, the profilers operated with five
beams. One beam is vertical and the others are disposed ev-
ery 90
◦in azimuth with a 77
◦elevation angle. The wind pro-
filers provided vertical profiles every five minutes of the three
wind velocity components, the reflectivity and the Doppler
spectral width, from 75 m above ground level (AGL) up to 2
or 3 km, depending on atmospheric conditions, with a 75-m
vertical resolution. The meteorological peak in the Doppler spectra is selected according to a 30-min duration consensus technique, mainly based on thresholds, vertical and temporal continuity. More details on the data processing, radar tech- niques and signal scattering can be found in Jacoby-Koaly et al. (2002) and Puygrenier et al. (2005). In addition to the basic measurements of the wind profilers, the turbulent ki- netic energy dissipation rate
εcan be deduced from the UHF Doppler spectral width and reflectivity is proportional to the air refractive index structure coefficient
Cn2. Local maxima of this coefficient are usually associated with the atmospheric boundary layer top,
Zi, resulting from high vertical gradi- ents of temperature and humidity and strong turbulent mix- ing. Then, the height-time section of the structure coefficient
Cn2enables one to follow the daytime ABL development with the accuracy of the
Zimeasurement being of the order of half the radar pulse length (75 m) (Angevine et al., 1994; Heo et al., 2003).
So, profilers that worked continuously during the whole campaign were an essential tool in the determination of the temporal variation of the ABL top, whereas radiosounding and aircraft vertical profiles could only provide sparse obser- vations.
2.2 Aircraft
To fulfill the ESCOMPTE objectives and to provide mea- surements on a regional scale, three airplanes equipped with instrumentation flew three times a day, following five hor- izontal and parallel legs along a west-east direction, at the constant level of 800 m (Fig. 1). The Fokker 27 ARAT from the Institut National des Sciences de l’Univers (INSU) flew at 05:00 and 11:00 UTC, whereas the Dornier 128 from the Institut f¨ur Meteorologie und Klimaforschung (IMK) flew in the early afternoon at 14:00 UTC. An exception occurred on J22, when the morning flights at 05:00 and 11:00 UTC were performed by the Merlin 4 from M´et´eo-France and the Dornier 128, respectively, and no flight was operated in the afternoon. The three airplanes measured classical meteo- rological parameters, radiative fluxes and trace gases (O
3, CO, NO
x)at a sampling rate of 1Hz. Turbulence measure- ments were recorded at a sampling frequency of 25 Hz on the Dornier and Merlin and 16 Hz on the Fokker. More de- tails about the aircraft instrumentation can be found in Sa¨ıd et al. (2005). These “exploration” flights allowed us to retrieve two-dimensional maps of chemical and meteorological vari- ables over the region, using a kriging method.
Spatial interpolation is rather tricky when dealing with me- teorological data measured over complex terrain. The krig- ing technique was used to provide fields of various param- eters measured on board aircraft, at the flight level. It is a so-called optimal interpolation method which predicts in- terpolated values from data observed at discrete locations (Matheron G., 1963). We used the Matlab kriging toolbox 3.0, developed by Lafleur and Gratton (1998) and almost en-
Fig. 1. Map of the various instrumented sites and horizontal track of the flights performed by the aircraft during the ESCOMPTE field experiment. Ellipses represent aircraft sounding locations. Trian- gles indicate radio-sounding sites and black half-moons correspond to the location of UHF wind profilers. Grey rectangles indicate the surface flux stations. Names of town, mountains and valleys are written in black, white and grey, respectively.
tirely written from functions proposed by Deutsch and Jour- nel (1992) and Marcotte (1991). The spatial length scale resolved by this interpolation method is about 12 km which corresponds to the distance between two parallel flight legs and thus allows a reasonable interpolation of the data in both directions (parallel and perpendicular to the flight legs). In addition, the standard deviation of the calculated fields yields information on the estimated error of the interpolated distri- bution pattern (Kalthoff et al., 2005). In our case, the stan- dard deviation for all fields is weaker than 10% and high- lights the reasonable accuracy of the method.
The turbulent kinetic energy (TKE) dissipation rate
εturned out to be useful for this study, since it helped to dis- criminate between the ABL and the free troposphere: turbu- lence is strong within the ABL and close to zero in the free troposphere. Temperature and humidity were used to find a threshold for
ε, using vertical profiles of the three parame-ters. The TKE dissipation rate
εcan vary over several orders of magnitude which is not the case in thermodynamics. In addition, a spatial increase in potential temperature (or de- crease in humidity) may be due to a horizontal gradient, as well as to a change in the vertical layer, so thermodynam- ics can hardly be used to discriminate between ABL and the free troposphere. We plotted fields of Log(ε) and observed that the spatial variability of TKE showed sharp transitions.
Then, combining these fields of TKE with the temperature and water vapor fields allows one to determine whether the aircraft flew within or above the ABL.
ε
was computed using a method described in Shaw
and Businger (1985) and Lambert and Durand (1999). It
consisted of band-pass filtering the time series of the vertical velocity component, in order to keep only the signal frequen- cies involved in the inertial subrange. The vertical velocity was filtered in the wavelength range 6 m to 100 m.
εwas then computed from this signal on short overlapping segments of 3 km, which corresponds to the model resolution.
2.3 Radiosoundings
Four sites were equipped with rawinsonde systems (Fig. 1).
Two rawinsonde systems were operated by IMK every 3 h from 08:00 to 20:00 UTC: one was located in the north- eastern part of the domain, at Vinon in the Durance val- ley, the second one was at Saint R´emy, close to the north- western ridge Les Alpilles. At Aix-les-Milles, in the central part of the ESCOMPTE domain, the Centre National de la Recherche Atmosph´erique (CNRM) also launched two ra- diosoundings per day between 06:00 and 17:00 UTC, with ozone profile abilities once a day. Furthermore, in the con- text of the Urban Boundary Layer (UBL) program (Mestayer et al., 2005), a rawinsonde system provided occasional mea- surements of temperature and moisture profiles at the Obser- vatoire in the Marseille city center.
2.4 Ground-based stations
We also used data from three ground-based stations which provided measurements of standard meteorological variables (air temperature, mean wind speed, humidity, etc. . . . ) and flux estimates every 30 min. One of the stations was in- stalled in Meyrargues (43.65
◦N, 5.5
◦E), in the Durance val- ley. Measurements were carried out above a maize field.
A complete description of the site can be found in Brut et al. (2004). The sites of Vallon d’Ol (43.35
◦N, 5.40
◦E), on Marseille foothills and La Crau, in the plain near Arles (43.57
◦N, 4.82
◦E) corresponded to bare soil. The instru- mentation implemented on these sites is described in Cros et al. (2004). These measurements sites are located in Fig. 1.
2.5 Complementary modelling
To further analyze study cases of IOP2, the three- dimensional mesoscale model Meso-NH was used. This model is a non-hydrostatic model, jointly developed by the Laboratoire d’A´erologie (CNRS) and the CNRM (M´et´eo- France) (Lafore et al., 1998). The chemical modules are coupled online with Meso-NH, which means that the meteo- rological and chemical fields (Suhre et al., 1998) are simul- taneously computed at each time step and each grid point. In our case, we used the gas-phase chemical module ReLACS (Crassier et al., 2000). The Genemis anthropogenic emis- sion inventory (Wickert et al., 1999) at 9-km resolution was used. Two-way grid nested simulations were carried out and provided two grids of resolution, 3 km and 9 km, centered on Marseille and Berre Pond. The domain of interest (the smaller) is located between latitudes 42.7
◦N and 44.3
◦N,
and longitudes 4.22
◦E and 6.42
◦E. There are 54 vertical lev- els, among which 40 lie between the surface and 2500 m. The simulation started at 00:00 UTC, on J21. The description of the model configuration for the simulation can be found in Cousin et al. (2005).
3 Description of IOP2
IOP2 (22 June 2001–26 June 2001) of the ESCOMPTE campaign is divided into two periods (hereafter IOP2a and IOP2b), with significant changes in the meteorological situ- ation on J23 around 17:00 UTC (Cros et al., 2004). IOP2a includes J22 and 23 and it corresponds to the end of a Mistral episode, i.e. moderate north-west wind. The IOP2b (24 June 2001–26 June 2001) corresponded to a different airflow situ- ation, dominated by the sea breeze which best characterized a pollution event. In this section, we briefly recall the mete- orological situations encountered during these five days.
The synoptic conditions observed from J22 to J23 con- sisted of a north-west regime at 500 hPa with a ridge from Gibraltar to Ireland and an extended 1013 hPa isobar over the Mediterranean Sea. This resulted in a case of Mistral, a northerly low-level wind induced by orography. This char- acteristic outbreak of south-eastern France often brings cold and dry continental air masses and restores clear sky condi- tions after the passage of a cold front. A complete descrip- tion of the Mistral events that occurred during ESCOMPTE, as well as its interaction with the sea breeze or its effect on air pollution transport can be found in Gu´enard et al. (2005), Bastin et al. (2006) and Corsmeier et al. (2005).
On J22, the weather was fair with a clear sky. Tempera- tures which were around 21 to 23
◦C in the morning, reached maxima of 31 to 33
◦C in the afternoon. The wind was weak and coming from the northeast at the beginning of the day.
In the early afternoon, a thermal breeze settled in the lower layer near the coast while the flow remained from the north- west inland.
On J23, the weather was still sunny and hot temperatures were observed with a maximum of 30 to 32
◦C over the whole region. In the morning, a weak north-westerly wind blew all day long over the region, except in the western part of the domain where it was stronger (5.1–7.7 m s
−1). In the after-noon, the sea breeze set up around 16:00 UTC and extended along the coast.
On IOP2b, western Europe was under the influence of a high pressure system centred over the Netherlands. Pres- sure gradients were weak over the whole of central Europe (Kalthoff et al., 2005). In fact, there was a complex interac- tion between synoptic northerly winds and local circulations;
these conditions favoured both the evolution of mesoscale circulation systems and the evolution of a summer smog sit- uation (Vogel et al., 1995).
On J24, the sky was clear from early morning: tempera-
tures were about 18
◦C in the western part of the domain and
15
◦C in the east. They increased up to 28
◦C and 30
◦C in the afternoon. Between 500 m and 1500 m we observed a mod- erate flow from the northwest (5.1 to 7.7 m s
−1)which weak- ened and veered from the west-northwest. In the lower lay- ers, the sea breeze set-up along the coast and then extended inland with a south-west direction and a moderate strength (5.1 to 7.7 m s
−1). This sea-breeze day was also describedand studied by Cousin et al. (2005).
On J25, an anticyclonic cell is located in the North Sea with an extension towards the Mediterranean basin: this in- duced a weak wind regime, dominated by a sea breeze. Since 07:00 UTC, a southerly wind has been observed near the coast, due to a sea-breeze effect; around 15:00 UTC, the sea- breeze has spread over the whole region. In addition, the sky remained clear with a few altocumulus clouds and tem- peratures from 26
◦C to 28
◦C near the coast and 31 to 34
◦C inland.
On J26, a weak pressure gradient at the surface induced a breeze regime. A southerly breeze rapidly replaced the weak to moderate morning wind from the east to the south- east, especially in the Rhˆone valley, whereas it was from the southwest in the eastern area. Eventually, the temperatures were about the same as the day before and a few altocumulus clouds were present in the sky.
Additional information and illustrations can be found in Bastin et al. (2005a) and Kalthoff et al. (2005) for J25 and J26, and in Delbarre et al. (2005) for J26.
4 Comparisons between modeling and experimental data at 2 levels
In this section, we compare the numerical simulations pro- vided by the MesoNH model with experimental data at two heights in the ABL (the surface and the 800 m level) on sev- eral days of IOP2.
4.1 Sea-breeze and Mistral on 22 and 23 June
The interactions between the sea breeze and Mistral have al- ready been studied in other publications (Bastin et al., 2006, 2005a, b; Caccia et al., 2005), whose authors focused mainly on J22.
Comparison between modelling and measurements at the surface and at the 800-m level
Time series of temperature, water vapor mixing ratio and wind speed and direction are displayed in Fig. 2, for J22 at the three ground sites (La Crau plain, Vallon d’Ol and Meyrargues). The distance of the three ground stations to the closest coastline is about 30, 10 and 40 km, respectively. The sea breeze is usually detected through a diurnal veering of the synoptic wind, which is accompanied by a temperature decrease and moisture increase. The onset times are repre- sented by an arrow on the figures and are summed up in Ta-
Fig. 2. Time series of meteorological parameters measured on 22 June above bare soil in the plain of La Crau (a and b), in Vallon d’Ol in the northern suburb of Marseille (c and d) and above a maize field in Meyrargues (e and f). On the left, the air temperature is represented with a solid black line and the water vapour mixing ratio with a dashed grey line. On the right-hand figures, the wind speed is represented with a solid black line, whereas the wind direction is represented with grey squares. The black arrows indicate the sea- breeze onset.
ble 1. So, the sea breeze reaches La Crau around 14:00 UTC and persists until 19:00 UTC. The sea breeze reaches Vallon d’Ol around 14:30 UTC when the wind and the temperature decrease while the humidity rate increases. This slight de- lay between the two sites is consistent with the observations of Bastin et al. (2006) but is not linked to the distance to the coastline. No sea breeze is detected in Meyrargues: the wind is from the northwest and corresponds to the Mistral.
The potential temperature diurnal cycle is bell-shaped, like
a classical surface layer over such a ground cover, whereas
it is sharper in the two previous cases. The humidity cycles
are very different over the three sites: in La Crau, we ob-
serve the usual humidity increase due to the sea-breeze on-
set. In Vallon d’Ol, the humidity decreases between 09:00
and 15:00 UTC and increases again with the sea-breeze on-
set. In Meyrargues, the humidity variability is also related
Table 1. Sea-breeze onset time (UTC) detected at the three ground stations/with the model for the whole period.
Ground station 22 June 23 June 24 June 25 June 26 June La Crau 14:00/14:00 15:30/– 09:00/09:00 07:00/09:00 07:30/07:30 Vallon D’ol 14:30/15:00 11:00/– 07:00/09:00 08:30/09:00 07:00/08:00 Meyrargues –/– 17:00/– 12:00/12:00 10:00/12:00 13:30/13:00
Fig. 3. Maps of wind speed and direction (m s−1), potential temperature (◦C) and water mixing ratio (g kg−1) at the surface (top) and at a 800-m height (bottom) simulated on 22 June at 12:00 UTC with Meso-NH. The quadrilatere delimits the area flown over by the aircraft.
to the maize transpiration and water deposition (for instance, the peak around 06:00 UTC is attributed to the morning dew).
These results are in good agreement with those from Bastin et al. (2006), who investigated a wider domain than ours.
They showed that the presence of the Mistral and other var- ious mechanisms that drive its dynamics in different regions of the Rhˆone valley lead to a very inhomogeneous and un- steady behavior of the sea breeze at the scale of the Rhˆone valley exit. They observed that on J22, the sea breeze set up around 11:00 UTC at the west of the Rhˆone valley exit (at the south of the Massif Central) and then penetrated later in- land, to the east of Marseille (14:00 UTC). They explained
this by a strong sea-land temperature gradient on the western area where the ground could have warmed up due to the wind sheltering effect caused by the mountains of Massif Central.
Figure 3 presents the thermo-dynamical fields provided by the MesoNH simulation on J22, 12:00 UTC, at the surface and at the 800-m level (in the following section, we will show that the 800-m level is within the ABL at 12:00 and 14:00 UTC for both days). Since no flight was operated on J22 at 14:00 UTC, we instead display the numerical fields at 12:00 UTC, so that we can compare them to the aircraft data which were recorded between 11:00 and 12:30 UTC.
At the surface, the model simulates a south-westerly wind
Fig. 4. Fields of wind speed (a), potential temperature (b) and water vapour mixing ratio (c) retrieved from aircraft measurements at 800 m on 22 June between 11:00 UTC and 12:30 UTC. The flight track is represented with a black solid line.
penetrating inland in the western part of the coast, to the west of the Berre Pond (Fig. 3a). The wind is from the west at the Marseille coastline. Inland, there is a very weak northerly flow in the Rhˆone valley which strengthens and turns north- west over the hills and mountains. In addition, the poten- tial temperature field inland exhibits a west-east gradient that matches the relief (colder in the Rhˆone valley than over the mountains) (Fig. 3b).
At 800 m, the dynamic and thermodynamic fields are quite different: the wind (Fig. 3d) is northerly to north- westerly with a magnitude of 6–7 m s
−1in the west, and up to 12 m s
−1in the east of the domain. The temperature field (Fig. 3e) displays a channel of a minimal value (28
◦C) in the Rhˆone valley and maximum values in the east (31–32
◦C):
the west-east gradient mentioned at the surface between the plain and the mountains is even more marked. At this level,
the temperature gradient is due to the differential heating be- tween the Rhˆone valley and the mountains (Alps to the east).
This gradient vanishes over the sea and the flow corresponds to the Mistral exit out of the Rhˆone valley. So, the sea breeze does not affect the atmospheric layer at this altitude and this is confirmed with the humidity parameter (Fig. 3f): the field of water vapor mixing ratio is indeed rather homogeneous on the eastern region (both inland and over the sea), with a value of 6 g kg
−1and higher values (7–8 g kg
−1)in the Rhˆone val- ley and downwind over the sea. The comparison of potential temperatures between both levels (2
◦C warmer at the ground than at 800 m over the Rhˆone valley; 3
◦C east of Vinon) sug- gests a strong instability in the ABL which is likely to yield high sensible heat fluxes inland.
The simulated fields at 800 m compare well to the air-
craft data. On Fig. 4, the measured wind is blowing from
Fig. 5. Same as Fig. 2 for 23 June.
the northwest with a magnitude of 5–6 m s
−1at the exit of the Rhˆone valley and a magnitude of about 12 m s
−1in the central part of the mountains, showing good agreement with the numerical results. The potential temperature presents a west-east gradient, with temperatures increasing from 29
◦C to 31
◦C, above the massifs of Sainte Baume and Sainte Vic- toire. Eventually, again in accordance with the numerical results, the humidity field is rather homogeneous, with the value of the water vapor mixing ratio varying around 6–
7 g kg
−1, slightly higher at the exit of the Rhˆone valley.
In the early afternoon (14:00 UTC, not shown here), the model can simulate the sea breeze at the surface but it is restricted to a short band along the coast between Saint Chamas and the east of Marseille. In fact, the onset occurs at 14:00 UTC in La Crau and 15:00 UTC in Vallon d’Ol (Ta- ble 1). As expected, the sea breeze is not observed in the numerical fields at 800 m, at 14:00 UTC.
On J23, the meteorological situation is rather similar to J22. Figure 5 displays the time series of wind, temperature and humidity at the same ground sites on J23. The sea breeze clearly appears in La Crau around 15:30 UTC with a change in the wind direction from the north to the west, a sharp in- crease in the water vapor mixing ratio associated with a de-
crease in potential temperature. The wind is veering again around 19:00 UTC, with a south-westerly direction. The air mass is even colder and moister. In the foothills of Marseille, the sea-breeze is detected at 11:00 UTC with an increase in the humidity rate associated with a decrease in temperature.
The wind blows from the southwest until 20:00 UTC with no change in the humidity rate. The sea breeze reaches Meyrar- gues at 17:00 UTC: the wind starts veering from the north- west to the south, and the temperature decreases while the humidity increases. Although the sea breeze sets up earlier in Marseille on J23 compared to J22, it penetrates later in the Rhˆone valley and it even reaches the Durance valley en- trance.
Since the sea-breeze onset occurred later inland than on the previous day, the 14:00 UTC situation is presented here.
The Meso-NH simulations at 14:00 UTC do not show any sea-breeze penetration at the surface (Fig. 6), whereas it slightly penetrated inland on J22. The Mistral is well estab- lished in the Rhˆone valley channel with a speed of 6–7 m s
−1. In Fig. 6b we can notice that the Mistral flow warms up as it blows above the warm ground of the Rhˆone valley exit. Therefore, this induces a large gradient of temperature and humidity when the Mistral air mass reaches the coast and meets the sea-breeze circulation, which conveys wetter and colder air. This gradient is associated with a west-east temperature (Fig. 6b) and humidity (Fig. 6c) gradient be- tween the Rhˆone valley and the eastern mountains, even more marked than on the previous day.
At 800 m, on J23, the numerical simulations show that the air mass from the Rhˆone valley is advected over the marine atmospheric surface layer. For instance, in Fig. 6d, the wind is orientated from the north to the northwest with an acceler- ation in a narrow band between Avignon and Marseille (11–
12 m s
−1), whereas its speed reaches 8 to 9 m s−1at the bor- ders of the domain. Compared to the aircraft measurements, a major difference appears in the wind field (Fig. 7a). Al- though the direction and the structure (i.e. the channelling effect along the axis “Saint Remy-Marseille”) are correctly reproduced, the measurements provide lower values of the wind speed (between 5 to 9 m s
−1)compared to those simu- lated (6–13 m s
−1). At 800 m, the temperature field obtainedwith the model (Fig. 6e) shows that the air mass has a mini- mum of potential temperature in the western part of the over- flown area (31
◦C) and a maximum in the eastern part (34
◦C).
This west-east gradient is clearly confirmed by the aircraft
data field (Fig. 7b). On the humidity field (Fig. 6f), we find
an east-west gradient (like the day before) with the mixing
ratio of water vapour reaching 7 g kg
−1in the Rhˆone val-
ley and 3–4 g kg
−1to the east. However, this gradient is not
observed on the aircraft measurements (Fig. 7c), which do
not show any organization but display humidity values in the
same range (3–7 g kg
−1).Fig. 6. Same as Fig. 3 for 23 June, 14:00 UTC.
4.2 The sea-breeze front along the coast during IOP2a To further analyze the differences in the sea-breeze onsets on J22 and J23, we investigated the sea-breeze vertical struc- ture. Figure 8b displays the height-time cross section of the horizontal wind measured with the UHF profiler in Saint Chamas. The sea-breeze westerly winds can be observed from 15:00 UTC on J22 in a layer up to 500 m and around 17:00 UTC on J23 in a thinner layer (up to 400 m). So the sea-breeze appears at around the same time in Saint Chamas (15:00 UTC), La Crau (14:00 UTC) and Vallon D’ol (14:30 UTC) on J22, whereas the onset times are 17:00, 15:30 and 11:00 on J23. Figure 9 shows a height-latitude south-north cross section of the simulated horizontal wind along an axis Marseille-Aix en Provence on J22 (a) and J23 (b) at 15:00 UTC (the horizontal track of the cross section is shown in Fig. 16). The colored scale represents the po- tential temperature. On 22 June, the sea-breeze appears as a westerly wind layer, with a stratified potential temperature structure (blue color) which extends up to 200 m in altitude and to the northern foothills of Marseille in latitude. It does not penetrate far inland and does not reach Aix en Provence.
The atmosphere is stratified over this marine layer and the
wind is a composition between the Mistral (blowing from the north-west at 1000 m) and the westerly sea-breeze flow.
Over the continent, a mixed boundary layer is developed in the Mistral flow from 1200 m (north to Marseille) to 1800 m, at the northernest point of the section.
On 23 June (Fig. 9b), the simulation shows that the sea breeze has hardly penetrated in Marseille, although it has been detected since 11:00 UTC with in-situ measurements (at Vallon d’Ol). It remains confined in the very lowest levels of the atmosphere. This effect is enhanced in the simulation since the Mistral magnitude is overestimated by Meso-NH, as noted in the previous section. The stratified layer capping the sea breeze layer is disorganized by the continental layer that is advected over the sea by the mistral flow.
Thus, for these two days, the sea breeze occurs during
the afternoon (except in Marseille on J23), roughly at the
same time at the scale of the studied area. No real sea-breeze
layer develops because the sea breeze is not strong enough to
struggle with the Mistral. However, the marine surface layer
that penetrates over the continent disconnects the boundary
layer from the surface. The Mistral is also advected offshore,
over the sea-breeze layer, on J23.
Fig. 7. Same as Fig. 4 for 23 June, 14:00 UTC.
4.3 Validation of the model on a pure sea-breeze case: 25 June 2001
The meteorological situation changed during IOP2b and turned to well-marked sea-breeze conditions. As indicated in Table 1, the sea breeze beginning is detected earlier on J24 than on the previous days (around 09:00, 07:00 and 12:00 UTC at La Crau, Vallon D’Ol and Meyrargues, re- spectively). On J25, which is a pure sea-breeze event, and J26, which corresponds to a synoptic situation from the south, the sea breeze starts early in Vallon d’Ol and La Crau and reaches Meyrargues more or less later in the morning.
Figure 10 displays the meteorological parameters measured at the ground-based stations in La Crau, Vallon d’Ol and Meyrargues on J25. The temperature cycles in Meyrargues and La Crau are bell-shaped, contrary to the rather flat one observed in Marseille. Actually, we can observe this flat type of cycle during the whole IOP2b: the sea breeze in Marseille
prevents the potential temperature from increasing. On the Meso-NH temperature fields at the surface (Fig. 11b), we can note that the temperature gradient is very strong and homo- geneous along the coast. Therefore, the advection of colder maritime air penetrates further inland compared to IOP2a.
The field is homogeneous inland with values around 34
◦C at 14:00 UTC.
The sea-breeze direction is from the west (i.e. perpendic- ular to the coast) in Marseille and from the south in La Crau and southwest in Meyrargues. The simulations at 12:00 UTC and 14:00 UTC show similar results as the measurements.
The ground measurements of Vallon D’ol are representa-
tive of the sea-breeze flow in the eastern part of the domain
(south-west direction) but are slightly different from the ob-
servations and model results (westerly flow) downtown in
Marseille, due to the direction of the coastline (Delbarre et
al., 2005).
Fig. 8. Height-time cross sections of the structure constant Cn2(m−2/3) (a), horizontal wind (m s−1) (b), vertical velocity (m s−1) (c) and dissipation rate of turbulent kinetic energy (10−4m2s−3) measured with the UHF wind profiler in Saint Chamas during the Mistral episode (J21 to J23). The red solid line in panel (a) represents the ABL top height.
In general, due to the weak values of the wind speed (2 to 5 m s
−1), there is a strong spatial variability in the winddirection on J25. At 14:00 UTC, the wind is very weak in the northern part of the domain (Fig. 11a–d) that is not reached by the sea breeze yet. In the eastern part, a well marked south flow generates a convergence area over the eastern relief. In the southwest part of the area the sea breeze can reach 5 to 7 m s
−1.
Contrary to the situation of IOP2a, we do not observe a strong increase in the humidity rate with the sea-breeze onset, especially in La Crau plain (Fig. 10a). In Meyrargues, be- cause of the transpiration of the maize, the water vapour mix- ing ratio is higher than at the other sites. These data compare well with the numerical fields of temperature and humidity at the surface (Fig. 11b–c). On the modelling results, we find again a strong humidity gradient along the coast but there is no noticeable advection of moist air due to the sea breeze.
The mountainous areas in the east seem to be moister than in the region north of the Berre pond (between Saint Chamas and Aix en Provence). At the 800-m level, the humidity rate is higher over the continent, which is not surprising consid- ering that the air masses are completely disconnected: as we will see in the next section, 800 m is either inside or over the internal boundary layer generated by the sea breeze.
It is more difficult to correctly simulate the wind field dur- ing IOP2b, because the wind has weakened and become un- steady. The sea breeze sets up earlier than during IOP2a but
local winds are dominating and the vertical structure of the low troposphere is very heterogeneous, even inside the sea- breeze layer.
In the next section, we will show that the decoupling be- tween the surface and the 800-m level is essentially due to the various stages of sea-breeze penetration.
5 ABL development
5.1 ABL height: definition and estimation
This section is devoted to the description and the analysis
of the ABL development over the ESCOMPTE area during
IOP2a and IOP2b. The depth of the ABL is a crucial pa-
rameter since it drives the volume of air within which the
pollutants are likely to be either diluted or concentrated. It
is indeed a hard task to define a ABL height in Marseille’s
surroundings where you can find together a big city, a com-
plex coastline and hilly terrains. Delbarre et al. (2005) high-
lighted the complexity of the air circulation in this area un-
der sea-breeze conditions. By studying the extinction of a
UV lidar signal along with the turbulence and reflectivity
measurements of a UHF radar and radiosounding data, they
showed, on J26, the existence of a 300-m deep layer with
a westerly sea breeze, capped by a southerly flow extend-
ing to 800 m or 1000 m above the ground level. Those two
Fig. 9. Vertical section of potential temperature (color scale) and horizontal wind simulated by Meso-NH along an axis Marseille- Aix en Provence on 22 June at 15:00 UTC (a) and on 23 June at 15:00 UTC (b). The location of the cross section is shown in Fig. 16b.
layers were characterized by different turbulence properties and were separated by thin stable layers, inhibiting the ex- changes. Lemonsu et al. (2006) also highlighted this super- position of two sea breezes, using a large eddy simulation with a 250-m resolution in the vicinity of Marseille.
Therefore, the difficulty of the present work was to find a means of defining an ABL in such complex situations, or at least a level likely to indicate a disconnection between an area of local exchanges (exchanges between the various internal layers) and the free atmosphere. Since we did not
Fig. 10. Same as Fig. 2 for 25 June.
observe conditions of a classical mixed layer for this defini- tion (which would mean a constant potential temperature and constant scalars concentration), we define
Zi(which usually corresponds to the ABL top) as the capping level of all the turbulent superimposed layers. This level corresponds to a sharp inversion that clearly indicates the separation with the stable layers of the free troposphere. This multi-layer ABL topped by level
Zican be considered as a reservoir whose chemical species may exchange from one layer to another.
For the whole IOP2, the ABL height
Ziwas determined using both numerical and experimental results. We used profiles of potential temperature and humidity from radio- balloon- and aircraft-soundings (Fig. 12), as well as time- height cross sections of reflectivity provided by the UHF radars (Fig. 8a). In addition to the available measured data,
Ziwas estimated using the vertical sections of potential tem- perature provided by the Meso-NH simulations. From these vertical sections, the ABL height is obtained from the max- imum of the temperature gradient which usually follows the
“ridge” of the isothermal lines along the section. However, it
remained difficult to define standard criteria to determine the
ABL depth in an area where it varies very sharply.
Fig. 11. Same as Fig. 3 for 25 June, 14:00 UTC.
Nevertheless, to retrieve maps of the ABL depth from modeling results, we used zonal and meridian vertical sec- tions of the potential temperature every 0.1 degree in longi- tude and latitude, respectively. The values of
Zi−Zr(where
Zris the ground surface height) were linearly interpolated over the ESCOMPTE domain, in order to provide maps of the ABL depths at 11:00 and 14:00 UTC for IOP2a and IOP2b. These fields are shown, respectively, in Figs. 13 and 14. This representation in terms of depth of the ABL instead of height has been preferred to outline the effect of the com- plex terrain in the following. The experimental results are also reported on these maps.
5.2 Comparison between model and experimental results The simulated ABL depth fields were validated using the ex- perimental data, at 11:00 and 14:00 UTC, which coincides to the times at mid-flight. The comparison is shown as bar diagrams in Fig. 15 and the difference is reported in Table 2, in terms of a mean relative difference between simulated and experimental data. A distinction is made between the results in the plain and near the coast and the results obtained in Vi- non, in a mountainous surrounding. The bar diagrams are
Table 2. Mean relative error between simulated Ziand measured Zi using the experimental data collected at St R´emy, Arles, St Chamas, Marignane and Aix-Les-Milles (without Vinon) and in the mountain area (Vinon).
1Zi(exp-sim)/Ziexp 22/06 23/06 24/06 25/06 26/06
11:00 UTC 0.12 0.03 0.22 0.27 0.11
(without Vinon)
11:00 UTC 0.10 0.00 0.91 0.33 0.00
(at Vinon)
14:00 UTC 0.22 0.09 0.08 0.05 0.00
(without Vinon)
14:00 UTC 0.12 0.40 0.45 0.43 0.08
(at Vinon)
presented for seven locations, which are indicated according to their name, their longitude or their shortest distance to the sea.
At 11:00 UTC or 14:00 UTC, the difference between
model and experience does not exceed 20% for 60% of the
results (Table 2). The ABL height is usually well estimated
Fig. 12. Potential temperature profiles from radio-soundings launched at Saint R´emy (top) and Vinon (bottom) during IOP 2a (left): J22 (solid black line) and J23 (dashed blue line) and on J25 for IOP 2b (right) at 06:00, 09:00, 12:00 and 15:00 UTC.
inland over the plain (Avignon, Arles, St R´emy) (left part of the bar diagrams in Fig. 15), whereas the largest errors are often due to an underestimation of the ABL near the coast at 11:00 UTC (central part of the bar diagrams). This is not surprising during IOP2b, since the sea breeze is in a transi- tion phase at this hour of the day. At 11:00 UTC the sim- ulated sea-breeze front may not be exactly positioned at the right location, since the simulated onset time is sometimes delayed in comparison to the observations (Table 1: Vallon d’Ol on J24 and J26, for instance). This argument is con- firmed by the results at 14:00 UTC which show a real im- provement in the ABL estimation near the coast. Actually, the error range, which also includes some of the inland sites, such as St R´emy, Arles or Avignon, decreased from 3–27%
to 0–9% from 11:00 UTC to 14:00 UTC. However, it must be noted that the ABL estimates at 14:00 UTC for J22 are re- moved from these results, as they present large discrepancies due to the Marseille case. This will be discussed below.
The case of Vinon is shown apart due to the specific be- haviour of the ABL over the mountain. Both model and ex- perience display an obvious west-east positive gradient in
Zi, (see, for instance, Fig. 15 at 11:00 UTC). This is not surpris- ing, owing to the increase of the relief between the Rhˆone valley and the eastern mountain. We will see in the follow- ing that this increase in the
Zi(ASL) level is not only due to the relief but also to the deepening of the ABL over the mountain. However, the model tends to overestimate this in- crease at 11:00 UTC as at 14:00 UTC, especially on J24 and
J25, with an error often reaching 40% (and even 91% once) (Table 2). This situation is due to some specific features of the simulated wind fields for these days: they exhibit a con- vergence area on the eastern mountains between the westerly (J24) or south-westerly (J25) sea-breeze flow and a south or south-east flow, to the east of the domain. The observations do not confirm this convergence area. On J24, the observed field exhibits a west-east gradient consistent with the ups- lope wind generated by the temperature gradient between the Rhˆone valley and the mountains. This is favourable for an in- crease in the ABL over the relief, but this increase is not as strong as the one induced by the simulated convergence.
To further validate the spatial variability of the ABL height
retrieved from the simulations, we also used fields of turbu-
lent kinetic energy (TKE) dissipation rate, estimated from
aircraft data at the height of 800 m. This parameter does
not provide the ABL depth but it is a means to discriminate
whether the 800-m level (flight level) is inside or outside the
ABL. This was very useful near the seashore where very few
vertical sounding were performed and where the ABL was
often very low due to the sea breeze. An example of the field
of the TKE dissipation rate obtained with the aircraft on J25
around 11:00 UTC is shown in Fig. 16a. It clearly shows that
the ABL is higher above the mountainous area, to the east to
St Chamas and over a narrow strip in the west correspond-
ing to the Alpilles ridge. To aid the comparison with the
Zi-height retrieved with the model (Fig. 16b), we added on
the latter the line corresponding to
Zi=800 m. There again,
the Alpilles ridge is shown to be entirely inside the ABL at 800 m, whereas Avignon or the coastal area south of the Alpilles are outside the ABL at this level. Aix-les-Milles is just at the border, both for the model and the experiment.
So, we can conclude that the model results and measure- ments are in good agreement on the whole and when dis- crepancies occur, this can be attributed to either a phase de- lay between the simulations and the measurements or to a wrong simulation of the flow around the mountains under weak wind conditions. This conclusion allows us to use the numerical results for our further comments on the ABL de- velopment.
5.3 Spatial and temporal variability of the ABL develop- ment
The previous comparison began to highlight different ABL heights according to the location relative to the coast or to the relief, according to the meteorological situation (IOP2a or IOP2b) and finally, according to the time of day and the various diurnal evolutions of the ABL. The following study relies upon the radiosoundings in Fig. 12 (ABL heights) and the maps of
Zi−Zr(ABL depths) presented in Figs. 13 and 14 for both IOPs. Several striking features can be pointed out:
5.3.1 IOP2a – different ABL behaviors over the plain for the two days
Although the meteorological conditions are very similar on J22 and J23, the ABL developed differently over the plain.
The ABL depth hardly reaches 600 m in the westerly part of the domain on J23 (Fig. 13), whereas it could reach 1500 m at the same hour in Arles the previous day. The first day of IOP2a shows a typical development of a convective ABL un- der summer clear conditions. In Saint R´emy,
Ziincreases from 350 m at 06:00 UTC to 1600 m at 12:00 UTC (Fig. 12).
In Fig. 8, which shows the sketch of radar parameters mea- sured at Saint Chamas (but it is nearly the same in the other radar locations) for three successive Mistral days, the striking feature is the repetitiveness of some cycles in the horizontal wind organization, in the ABL height (at least for the first two days) and in the vertical velocity during the three days.
For instance, the ABL development can be followed using the time-height cross section of reflectivity (Fig. 8a). In this figure, the top of the ABL is deduced from the maximum of the reflectivity signal that reflects the sharp humidity gradi- ent at the inversion level. It clearly appears that the ABL in- creases faster and higher in Saint Chamas on 22 June than on 23 June. The rapid and deep development of
Zion J21 and J22 is due to strong, sensible heat fluxes (in Marseille, we ob- serve a maximum value of 450 Wm
−2on J22), resulting from the high difference in temperature and the specific humidity between the surface and the colder and drier air masses trans- ported by the Mistral wind. TKE is strong as well, as seen
in Fig. 8d. On J23, the ABL height shown in Fig. 12, or its depth presented on the left part of Fig. 13, remains very low over the plain, although the Mistral conditions persist with strong northerly winds (Fig. 8b) (the maximum value of the sensible heat flux in Marseille is 300 Wm
−2on J23). Ac- cording to Gheusi et al. (2004), the sea breeze has set up very early further to the west of the ESCOMPTE domain on J23 in an area where the Mistral wind has weakened due to the shel- tering effect of the Massif Central. Nevertheless, this is not the case in our studied area, except very close to the surface, as was previously shown with the ground stations (Table 1).
Even close to the coast, as in Marseille, the sea-breeze ver- tical extent, which could prevent the ABL from developing, does not reach 500 m before 15:00 UTC (westerly direction in Fig. 8b). However, in spite of persistent Mistral condi- tions in St R´emy, the ABL depth remains constant between 09:00 UTC and 12:00 UTC (Fig. 12). In fact, a strong in- version at 09:00 UTC inhibits the development of the ABL.
This strong inversion is linked to both a warming of the up- per layers (Mistral weakening in altitude) and a cooling of the surface conditions (Fig. 12: 29
◦C at 1500 m, 12:00 UTC on J22 for 33
◦C on J23; 22
◦C below 300 m, 06:00 UTC on J22 for 21
◦C on J23). In addition, a strong negative verti- cal velocity affecting the first 2 km of the low troposphere is maintained all morning long on J23 (Fig. 8d), whereas it disappears after 08:00 UTC for J22.
5.3.2 IOP2a – high spatial variability of the ABL for both days
If we now focus on the spatial variability of the ABL devel- opments for these two days, we can notice that the develop- ment is quite homogeneous on J22, with ABL depths ranging between 1000 and 1500 m and a variation that can be linked to the wind velocity distribution (Fig. 4): the ABL is the low- est when the wind is the strongest (12 m s
−1), showing thatdynamics prevails over thermodynamics.
The north-south oriented part of the Durance valley (in-
cluding Vinon) clearly displays some higher ABL, which are
favored by the constriction of the flow in the valley that ac-
celerates the wind and generates turbulence (Bastin et al.,
2005a). Another source of prevalent development is the
mechanism of valley breeze, which can be clearly seen on
the Vinon ABL heights shown in Fig. 12c.
Ziis very low at
09:00 UTC, due to the down-slope wind that brings cold air
to the valley bottom, inhibiting the ABL development. Be-
tween 09:00 UTC and 12:00 UTC, the differential heating
between the slopes and valley bottom, reverses the surface
flow and the ABL can grow rapidly, then it decreases again in
the early afternoon associated with the up-slope wind weak-
ening. This is supported by the surface wind in Meyrargues,
at the Durance valley entrance, which turns from the east to
the west on J22 between 06:00 and 09:00 UTC (Fig. 2f). It
starts decreasing from 14:00 UTC.
Fig. 13. Horizontal fields of atmospheric boundary layer depth obtained from the Meso NH simulations (colored scale) or experimental data (black labels), during IOP2a at 11:00 UTC (left) and 14:00 UTC (right).
Due to the low development of the ABL over the plain on J23 at 11:00 UTC, the difference in the ABL depths between the plain (500–600 m) and the mountain (1000 to 1400 m) is even more marked. The ABL in the Durance valley, near Vinon, remains higher than in the surroundings.
Another huge difference appears between both days, when looking at the 14:00 UTC maps (right part of Fig. 13): the ABL has highly deepened during the 3-h lapse time, reach- ing 2000 m in some areas. It preferentially develops over the plain on J22, and over the mountain on J23. On J22, it is due to the important heating of the air over the plain: the ABL depth has reached the temperature of the free troposphere, so that it can develop more easily (no inversion). On J23,
the ABL has hardly been able to break the inversion between 11:00 and 14:00 UTC over the plain, so it cannot increase as much as the day before.
Another source of wind circulation may occur at
mesoscale, due to the difference in radiative heating be-
tween the Rhˆone valley and the Alps foothill: 1.5
◦C on J22,
11:00 UTC (Fig. 4b); 4.5
◦C on J23 at 11:00 UTC and 5
◦C
at 14:00 UTC (Fig. 7b). This temperature gradient previ-
ously underlined between both areas may generate an up-
slope wind, roughly west-east, which could favor the ABL
development over the relief. This mesoscale flow is, how-
ever, inhibited when the synoptic flow is strong, as it has
been observed on J22: in this case, the ABL is not as deep
Fig. 14. Same as Fig. 13 for IOP2b.
Fig. 15. Comparisons of the boundary layer height Zi (ASL) determined from the Meso NH model (black) and from experimental data (grey) at 11:00 UTC (left) and 14:00 UTC (right). Values of Zi in different places over the ESCOMPTE domain are presented for the whole IOP from J22 (a) to J26 (e). The various locations of the sites are indicated using their name, their longitude and their distance to the seashore.
over the mountain as over the plain. On J23, 14:00 UTC, the Mistral has weakened over the mountain (5–7 m s
−1in- stead of 8–12 m s
−1)and the temperature difference has in- creased between the Rhˆone valley and the mountain, which favors the upslope circulation: the ABL develops preferen- tially over the mountainous part, reaching 2000 m in several places (Fig. 13).
5.3.3 IOP2a – offshore advection of the ABL
Also noticeable is the offshore extension of the continental ABL over the sea during this IOP. Figure 13 displays the apparent different behaviors of the ABL close to the coast- line between both days: on J22, the continental ABL ex- tends over the sea at 11:00 UTC. This is obvious on Fig. 13:
the ABL is still high over the sea at the point A18, where
the aircraft vertical sounding indicates that the ABL reaches
1400 m. This is not reproduced by the model which sim- ulates a sea-breeze penetration up to latitude 43.45 deg N at 4.7 deg E. The sounding result is confirmed by the air- craft’s horizontal field of the TKE dissipation rate
εat 800 m (not shown here), which indicates strong turbulence at 800 m near the coast. So, instead of a sea-breeze penetration, there is an advection of the continental ABL over the sea which is capping the marine surface layer. The only place where the simulation accounts for this advection is in the Mar- seille bay at 11:00 UTC. At 14:00 UTC again, the model fails to represent it: for instance, the UHF radar indicates an ABL depth growing from 1500 m to 2000 m between 11:00 and 14:00 UTC in Marseille (Fig. 15; see also Fig. 8a for St Chamas), whereas the model simulates its decrease from 1000 to 500 m due to the sea-breeze penetration (see also Fig. 9a from Sect. 4.2). The model indicates a competition between the sea-breeze layer and Mistral (with no real devel- opment of any sea-breeze layer), whereas the observations show that Mistral is simply advected offshore, above the ma- rine surface layer.
On J23, model and experience agree to display this off- shore horizontal advection of the continental ABL (Figs. 13, 9b). In addition, the numerical field at 14:00 UTC displays a deepening of the ABL in the advected continental layer, supporting the results from Caccia et al. (2004), who showed that the Mistral is raised up by the sea breeze near the coast.
The important feature to keep in mind is that for both days, when the synoptic wind is counteracting the sea-breeze, the effect of the latter is to advect offshore and raise the conti- nental layer. A well mixed layer of pollutants is transported over the sea. This layer is disconnected from the surface.
5.3.4 IOP2b – combination of various mesoscale circula- tions and consequence on the sea-breeze penetration During IOP 2b, the synoptic flow is rather weak, which al- lows the sea-breeze onset. The latter is provoked by the high difference in horizontal temperature between the sea-surface and the continental surface. Puygrenier et al. (2005) pre- sented the characteristics of the sea breeze measured in Mar- seille, in relationship with the horizontal surface temperature gradient between the sea and the continent: they indicated differences of 8
◦C, 7.7
◦C and 8.6
◦C for J24, J25 and J26, respectively.
However, this temperature difference that could be linked to the breeze intensity in the case of a rectilinear coast and flat area, is not the prevalent parameter in the ESCOMPTE area. Another important parameter in this area of complex to- pography is the presence of up-slope breezes, induced by the temperature contrast between the Rhˆone valley and the Alps foothills. This circulation, which was described before at the scale of the Durance valley, also developed as a mesoscale circulation, linked to the strong west-east temperature gradi- ent measured over the area: at 11:00 UTC, the west-east tem- perature gradient measured by the aircraft at 800 m is 3.5
◦C,
3.5
◦C and 4
◦C for J24, 25 and 26, respectively. It is not significant at 14:00 UTC, since the 800-m level is no longer inside the ABL but this gradient still exists at lower levels.
As a consequence, the wind measured over the ES- COMPTE area is the composition of moderate synoptic wind, with a sea breeze and up-slope breeze, i.e.:
– the end of the Mistral episode (north-westerly 6–
8 m s
−1synoptic wind, according to the Dornier flight at 2800 m, rapidly decaying in the ABL), with a sea breeze (south component at mesoscale) and slope-wind (west component) on J24: the result is a west moderate wind in the area at 800 m, with a channeling effect (ac- celeration) in the west-east and south-north parts of the Durance valley at 11:00 UTC.
– the synoptic wind is still north-westerly 6–8 m s−1
and decays during the day, according to the Dornier flight and the radiosoundings. It is very weak within the ABL.
The sea breeze is blowing from the south, as indicated by Bastin et al. (2005a, b) over the sea, with an up-slope breeze (west) on J25 which results in a weak, disor- ganized west, south-west wind, with again a channel effect in the Durance valley at 11:00 and 14:00 UTC (Fig. 11d).
– the synoptic wind is blowing from the south, southwest
